Quaternary Science Reviews 28 (2009) 2564–2581
Contents lists available at ScienceDirect
Quaternary Science Reviews
journal homepage: www.elsevier.com/locate/quascirev
Evolution of Ganges–Brahmaputra western delta plain: clues from sedimentology
and carbon isotopes
A. Sarkar a, *, S. Sengupta a, J.M. McArthur b, P. Ravenscroft c, M.K. Bera a, Ravi Bhushan d,
A. Samanta a, S. Agrawal a
a
Department of Geology & Geophysics, Indian Institute of Technology, IIT Kharagpur, Kharagpur 721302, India
Department of Earth Sciences, University College London, London WC1E BT, UK
Department of Geography, University of Cambridge, Downing Place, Cambridge CB2 3EN, UK
d
Physical Research Laboratory, Ahmedabad 380 009, India
b
c
a r t i c l e i n f o
a b s t r a c t
Article history:
Received 10 December 2008
Received in revised form
15 May 2009
Accepted 20 May 2009
Sedimentology, carbon isotope and sequence stratigraphic analysis of subsurface sediments from
western part of Ganges–Brahmaputra (GB) delta plain shows that a Late Quaternary marine clay and
fluvial channel-overbank sediments of MIS 5 and 3 highstands are traceable below the Holocene strata.
During the Last Glacial Maximum (LGM) sea-level lowering of >100 m produced a regional unconformity
(type 1), represented by palaeosols and incised valley. C4 vegetation expanded on exposed lowstand
surface in an ambient dry glacial climate. At w9 ka transgression inundated the lowstand surface
pushing the coastline and mangrove front w100 km inland. Simultaneous intensification of monsoon
and very high sediment discharge (w4–8 times than modern) caused a rapid aggradation of both
floodplain and estuarine valley fill deposits between 8 and 7 ka. The Hoogli River remaining along its
present drainage possibly acted as the main conduit for transgression and sediment discharge that was
subsequently abandoned. C3 vegetation dominated the delta plain during this time. From 7 ka onward
progradation of delta plain started and continued till recent. This period experienced a mixed C3–C4
vegetation with localized mangroves in the mid-Holocene to dominant return of C4 vegetation in the late
Holocene period. The study indicates that while the initiation of western part of GB delta occurred at
least 1 ka earlier than the global mean delta formation age, the progradation started at w7 ka, at least
2 ka earlier than thought before. The terrestrial vegetation change was modulated by changes in
depositional environment, specific ecological niches and climate rather than pCO2.
Ó 2009 Elsevier Ltd. All rights reserved.
1. Introduction
The sedimentary record within river deltas provide unique
opportunities to study the complex interplay among and control of
different forcing mechanisms like eustasy, climate and tectonics on
the development of depositional sequences of various cycles and
magnitudes (Bhattacharya, 2006 and references therein). Over the
last decade extensive research have been carried out in the delta
formed by Ganges–Brahmaputra (GB) rivers (Fig. 1a) of India and
Bangladesh (together these two rivers has the highest sediment
discharge capacity in the world i.e. w109 ton/year; Coleman, 1969).
These studies revealed important information about evolution of
the GB delta namely, the changes in sediment budgets/dispersal
* Corresponding author.
E-mail address: [email protected] (A. Sarkar).
0277-3791/$ – see front matter Ó 2009 Elsevier Ltd. All rights reserved.
doi:10.1016/j.quascirev.2009.05.016
through time and space and development of the lower delta plain
and coastal zones during the late Quaternary period (Allison et al.,
1998; Goodbred and Kuehl, 1998; Goodbred and Kuehl, 1999,
2000a; Allison et al., 2003). During the Last Glacial Maximum (LGM
w18 ka) sea level was lower than the present mean sea level (msl)
by at least w100 m and sediment delivery was extremely low due
to reduced water discharge through GB as a result of weak south
west monsoon and increased north-east monsoon which was
essentially dry (Cullen, 1981; Sarkar et al., 1990; Wiedicke et al.,
1999). The unique features of the GB delta system are that its
Holocene sedimentation presumably started at w11 ka, predating
all the major deltas by at least 2–3 ka and its shoreline was relatively stable in spite of rapid early Holocene sea-level rise (Goodbred and Kuehl, 2000b). This eustatic rise back flooded the
lowstand surfaces (as much as w70 km inland; Allison et al., 2003)
formed during the LGM resulting in an expanded estuary (Allison
et al., 2003). The Holocene shorelines were essentially traced out by
A. Sarkar et al. / Quaternary Science Reviews 28 (2009) 2564–2581
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Fig. 1. (a) Regional map of Ganges–Brahmaputra delta and location of present study area (Barasat); also shown the earlier studied bore hole locations (solid dots) used for sequence
stratigraphic analysis. Note paucity of data in western (Indian) delta plain. The Hoogli River acted as main conduit of sediment supply (palaeo-Ganges, shown by dark shaded region)
during initial delta growth in Holocene subsequent to which progradation proceeded from west to east (for details see text). (b) Map of the study area showing bore hole locations
spread in a N–S transect.
assemblage of mangrove pollen and marine shell fragments in
radiocarbon dated subsurface sediments (Vishnu-Mittre and Gupta,
1972; Banerjee and Sen, 1987; Umitsu, 1993). Yet the increased
accommodation space, created by rising sea level, was rapidly filled
up by enormous sediment discharge (w>2 times than present)
caused by intensified early Holocene (11–7 ka) monsoon (Van
Campo, 1986; Sirocko et al., 1993) driven by the regional insolation
maximum at w10–9 ka (Cohmap, 1988; Prell and Kutzbach, 1992).
These studies indicate that, apart from the sea level, regional
climate change has equally important role for sequence development in river deltas, a conclusion also supported by non-linear
numerical models of Goodbred et al. (2003).
In spite of these voluminous works a comprehensive understanding about the entire delta system is lacking. A closer look at the
studied sections will indicate that majority of the subsurface data
come from the eastern (Bangladesh) part of the GB delta (Fig. 1a).
Together, these data suggested abandonment and eastward migration of the active Ganges distributary (Goodbred and Kuehl, 2000a)
and the late Holocene progradation of lower delta plain in 4–5
phases. The earliest progradation (at w5 ka) took place in the
western extreme of the delta around the so called ‘‘early Ganges’’,
remnants of which are represented by river Hoogli in India and other
minor distributaries of the Ganges (Allison et al., 2003). Such
a model requires extensive study of subsurface sediment packages in
the western GB delta not only along the E–W tract of lower delta
plain but also in an N–S transect. Most studies in the western delta
confined to a generalized description of facies and sediment thickness and their possible connection with tectonics in Bengal basin
(Hait et al., 1996a, b; Stanley and Hait, 2000). Other studies mostly
reported the fossil pollen assemblages during the Holocene in terms
of vegetation change and shifts in mangrove front (Gupta, 1981;
Barui et al., 1986; Sen and Banerjee, 1988; Hait and Behling, 2005).
We observed that despite this work, the sedimentary successions in
the western delta plain still awaits a detailed facies and palaeoenvironmental analysis. Also, no attempt has so far been made to
understand the basin filling history of the western delta plain in
a space-time framework using sequence stratigraphic insight. Of
special interest is the types, fluxes and burial of organic carbon in
this delta which possibly are strongly interlinked with the Himalayan erosion and global atmospheric pCO2 change during the last
20 ka (France-Lanord and Derry, 1994; Galy et al., 2007, 2008a, b).
Since organic matter processing in estuaries/deltas is largely
controlled by climate, vegetation type (viz. mangrove vs. marine
particulate organic matter) and sedimentary processes (Dittmar
et al., 2001; Middelburg and Herman, 2007), it is indeed interesting
to see how the carbon cycle responded to the eustatic and
monsoonal changes over this time period.
In this paper, we attempt a detail facies analysis through several
drilled cores in the western GB lower delta plain that represent
a period from pre-LGM to recent. We also use high resolution
carbon isotope chemostratigraphy in these cores to retrieve the
change in vegetation vis-à-vis the climate and sedimentation
during LGM-Holocene time. Using the available and new radiocarbon dates, published sea-level data along the Indian east coast,
and sedimentary logs along a N–S transect a sequence stratigraphic
framework is proposed for the western delta plain.
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2. Geologic setting and regional stratigraphy
Bounded by Precambrian crystalline rocks in the north and west
and Assam-Arakan Neogene fold belt in the east, the GB delta
represents an integrated late Quaternary sedimentation history
over a 105 km2 area (Fig. 1a). Several uplifted Pleistocene terraces
occur within and bordering its alluvial plain. The delta is fed by two
major rivers viz. Ganges and Brahmaputra which supply large
quantities of water (w6.5 1011 m3) and sediment (w1 billion ton;
Coleman, 1969, Hossain, 1992) discharge into the delta plains and
also to the Bay of Bengal (BOB). The maximum discharges (w80%)
occur between June and October when summer (SW) monsoon
precipitation is highest (Sengupta and Sarkar, 2006). Monsoon
induced flooding causes widespread fluvial sedimentation in the
flood and delta plains. In addition to the Ganges and Brahmaputra,
several smaller rivers in the west also contribute to water/sediment
discharge, albeit to a lesser extent. The major river that drains the
western delta plain is Hoogli which alone accounts for about
20–25% of water and w65 106 tons of sediment discharge in the
GB delta system (Mukhopadhyay et al., 2006). Although much
lower than the Ganges and Brahmaputra combined discharges,
these figures indicate that this river too exerted substantial control
on sediment dynamics during the Quaternary-Holocene period.
Towards the Bay of Bengal, in the fringe area of the delta (both in
India and Bangladesh), occur Sunderbans, the world’s largest
mangrove forest and swamps (Fig. 1a). The name Sunderban comes
from the presence of dominant tree species, Heritiera, locally
known as ‘‘sundari’’ in these mangrove forests.
Earlier studies on subsurface stratigraphy of GB delta showed
presence of an oxidized sediment layer (interpreted as palaeosol in
McArthur et al., 2008, and the present study; see later discussion)
formed during the LGM lowstand. The palaeosol has been recorded
across the GB delta e.g. in Brahmaputra (Zheng et al., 2004),
Meghna (Yount et al., 2005), Jamuna (BADC, 1992) and Hoogli
(McArthur et al., 2008) floodplains. In the eastern part of the delta
this is overlain by thick (up to w60 m) silty sediments deposited
between 11 and 6 ka during the major transgressive event (Umitsu,
1993; Goodbred and Kuehl, 2000a, b). However, in the western part
(Indian side) the thickness of the Holocene sediments have been
found to be much less (w15–20 m; Hait et al., 1996a, b; Goodbred
and Kuehl, 2000a, b; Stanley and Hait, 2000). Goodbred and Kuehl
(2000a) suggested three major Holocene stratigraphic units from
the upper delta plain viz. a lower mud unit (max. 25 m) of 10–7 ka
age rich in peats, mangrove woods and marine fossils, a middle
fluvial sand-silt unit and an upper silt-mud deposit (max. 15 m).
The top two units supposedly correspond to the progradational
facies following the early Holocene transgression. While this gives
a generalized picture for the GB delta system, the progradation
history, particularly in the western part, is poorly constrained.
Although initiation of subaqeous delta clinoform has been suggested at w 7.5 ka (Michels et al., 1998), it is not known exactly
when progradation started in different parts of the delta. Further,
both the palaeosol and upper progradational units are not continuously present throughout the region and often eroded by single or
amalgamated fluvial channel sands of various ages (Stollenwerk
et al., 2007). While mud deposits are often well dated, chronological constrains for the channel sands are poor.
3. Materials and methods
Eight bore holes were drilled at Barasat locality (22 44.43/N,
88 29.45/E; Indian/Bangladesh datum; Fig. 1a), 20 km northeast of
Kolkata and w100 km north of the BOB coast (Fig. 1b) by reversecirculation, percussion method (Ali, 2003). The bore holes are
spread in an N–S transect adjacent to three villages Joypur, Ardivok,
and Moyna (together termed as JAM; McArthur et al., 2004, 2008).
The area falls in the western part of the southern Bengal Basin
(north 24-Parganas District in southern West Bengal, India) and
covered by modern alluvium. The southerly flowing Hoogli River is
located w15 km west of the area. Lithologs were prepared from the
retrieved cores (SW 1 to AP; Fig. 1b) and facies types were identified. ‘‘Key surfaces’’ (Transgressive surface of erosion, unconformity
etc.) were identified where abrupt change in interpreted bathymetry were found across the surfaces and taken as marker for
correlating different studied sections as well as to identify major
paleogeographic shift in the depositional history that can be traced
both regionally and basin-wide. 14C dating was made on two select
samples of organic rich clay by synthesis of sample carbon to
benzene following liquid scintillation counting method of Bhushan
et al. (1994) and dates were converted to calendar dates before
present (BP) using the calibration of Stuiver et al. (1998). A set of
five 14C dates (in wood/peat) and 4 OSL dates (from sands),
obtained in these cores, have already been published (see Table 1,
McArthur et al., 2008). Supplementary data Table S1 provides the
GPS locations of the cores, recovered length, materials dated and
isotopically analysed as well as ages obtained. Also compiled in
Table S1 published data for cores from seven different locations
across the western delta plain. Few samples have been subjected to
laser diffractometry grain size analysis. Six cores (SW 1,4,5,6,7 and
DP) have been analysed for carbon isotopic (d13C) compositions of
bulk organic matter (sands excluded). For this about 50 mg of
decarbonated sediment sample was combusted in a Flash
elemental analyzer. The evolved CO2, purified through a moisture
trap, was measured for its isotopic compositions in a Delta Plus XP
continuous flow mass spectrometer at IIT, Kharagpur. Few samples
from core SW 4 were analysed for d15N compositions following
same protocol and converting samples to pure N2. Analytical
precision for d13C and d15N is w0.1&. Also measured were the
total organic carbon (TOC), total nitrogen (TN) and ratio of total
carbon to nitrogen (C/N). For this samples were converted to CO2
and N2 in the elemental analyzer. The percentage of N and C were
calculated from the peak areas obtained from the sum of the m/z 28
and 29 or 44, 45 and 46 respectively measured in the mass spectrometer (Jensen, 1991). Typical analytical error was <1%.
4. Chronology and sedimentation rate
A total of 11 dates were available on the cores used from the
present study area (Table S1). Out of these four are OSL dates from
channel sands. The age of sediments ranges from w23 ka to w1 ka
spanning pre-LGM to latest Holocene. Because channel sands might
represent incomplete depositional record, the Holocene sedimentation rates were determined from overbank silt/clay deposits
wherever apparent continuity of sedimentation was observed. Two
cores DP and SW 4 have by far the thickest Holocene overbank silt/
clay deposits without any erosional channel sand (for lithologs see
Fig. 3) and have been used for this purpose. In order to ensure that
the sedimentation rate is indeed a regional representative, we have
also used eight available 14C dates from the nearest site Kolkata
(Table S1) where the vertical facies variation and overbank facies
thickness are comparable to those of Barasat. 14C ages are plotted
against depth in Fig. 2. The best fit lines through the data points,
taking into account the errors in 14C ages, show a distinct break in
sedimentation rate at 8 m corresponding to about 5.8 ka. The
calculated sedimentation rate in the mid- to late Holocene (w5.8–
1 ka) is w0.5 m/ka while the sedimentation rate in early to midHolocene (8–5.8 ka) is w4.4 m/ka, a large increase by a factor of
w9. The young age up to 4.5 m in Fig. 2 possibly indicates mixing of
topmost sediment layer by anthropogenic activity like modern
cultivation. Such mixing does not penetrate deep and preserved the
A. Sarkar et al. / Quaternary Science Reviews 28 (2009) 2564–2581
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5.1. FA1: Pleistocene shelf
Constituted entirely of gray to dark gray stiff clay this facies
(facies type A) represents the oldest stratigraphic unit in the area.
Although not shown in Fig. 3, the total thickness of this unit is
w30 m (obtained by deep coring elsewhere) and form the base of
all sections. The facies is laterally continuous, top eroded either by
sands of facies B (northern part of the area) or facies D in the
southern part. On the basis of well correlated TOC and total sulfur
(TS) defining a TOC/TS slope of w1.1, McArthur et al. (2008) interpreted this facies as marine clay. Striking lateral continuity and
fairly high TOC (w1.6%) suggests its deposition under dysoxic to
anoxic marine shelf (pro-deltaic clay? Kanellopoulos et al., 2006)
condition generally below storm wave base (Gawthorpe et al.,
2000). Upper part of this facies is oxidized with low TOC and TS
(Table S1) indicating exposure and weathering.
5.2. FA2: pre-LGM fluvial channel sand and overbank
Fig. 2. 14C ages of the Holocene sediments plotted against depth; data from both
Barasat and Kolkata are included. Note a factor of 4–8 increase in sedimentation rate
before 6 ka.
14
C chronology of sediment column older than 1 ka. Even ignoring
all the dates younger than 5 ka and joining the surface (‘0’ age) and
5.8 ka level (dashed line in Fig. 2) provides a much lesser (w1.2 m/
ka) sedimentation rate than the early Holocene time. In any case,
the estimated late Holocene sedimentation rates (minimum
w0.5 m/ka to maximum 1.2 m/ka) are quite similar to the 137Cs
based modern accretion rates (0.3–1 m/ka) from south-central
floodplains of the G–B delta plains obtained by Goodbred and Kuehl
(1998). Using the early Holocene rate of w4.4 m/ka the base of the
continuous Holocene clay sections at our area (core DP and SW 4
without any channel sand) is dated close to w9 ka. Since w9 ka is
obtained at the base of Holocene sections even at the most distal
locations (viz. Digha; Fig. 1a; Hait et al., 1996a) we consider this as
the initiation of Holocene sedimentation in this entire western
delta plain. Although near 10 ka date is found at the Holocene base
in eastern part of delta (Bangladesh; Umitsu, 1993; Goodbred and
Kuehl, 2000b), no date older than w9 ka has yet been obtained in
the western part. The significance of OSL dates on sands will be
discussed later in the context of facies characterization.
5. Facies and depositional environments
Supplementary Table S2 summarizes the facies types (classification scheme after Swift et al., 1991; Miall, 2000) and Fig. 3
illustrates facies dispositions in representative vertical lithologs
along a N–S (AP-SW 1; see Fig. 1b) section. Five major facies associations were recognized in the present study and each assigned to
specific depositional environment. The paleo-environments vary
from nonmarine alluvial plain to distal marine shelf viz. 1) FA1:
Pleistocene Shelf, 2) FA2: pre-LGM fluvial channel sand and overbank, 3) FA3: LGM palaeosol and incised valley fill, 4) FA4: Early
Holocene estuarine valley fill/aggrading fluvial channel sand,
overbanks and peat swamps 5) FA5: Middle to late Holocene
channel-floodplains with minor peat swamps. Due to the lack of
subsurface data only a single facies could be identified in FA1 and
was used to infer depositional environment.
Lenticular sediment bodies of this association (Table S1) overlies
FA1 sediments and are constituted of fine grained brown (ferruginous) sand with intervening layers of clay. The clay layers, however,
have not been encountered at the present study area but found at
other places like Kolaghat and Kolkata (see Fig. 7). Capped by facies
C, this is often laterally incised by w23 ka old sands of facies D
(Fig. 3). Low TOC content and ferrugination of the sand units
indicate deposition in an oxic environment. Chemically the sands
show much less concentrations in most major and trace elements
compared to younger sands of facies E (see later discussion;
McArthur et al., 2008) suggesting intense leaching. The clay layers
are often organic rich and contains abundant mangrove pollen like
Rhizopora, Avicennia, Heritiera etc. (e.g. at Kolaghat; Hait et al.,
1996b) indicating back swamps associated with a high water table
condition. At Kolaghat the clay layer of this facies association,
containing the above biota, is dated as w31.75 ka (Hait et al.,
1996b). The sand:mud ratio is high in the study area suggesting
possible amalgamation. At Kolaghat the ratio is relatively low (Table
S1) indicating isolated sand bodies. We infer a fluvial channeloverbank deposit for this association in general and low to high
amalgamated channel sands in particular, variously preserved at
different locations (Wright and Marriott, 1993).
5.3. FA3: LGM palaeosol and incised valley fill
This is composed of facies C and D. Facies C is represented by
brown clay with abundant decayed roots in its lower part (sub-facies
C1) with mildly calcretised top (sub-facies C2). It caps facies B and
eroded or overlain by facies E and F. Abundant decayed roots, low
TOC content (0.2–0.6%), relative enrichment in immobile elements
e.g. Fe,Cr,Ni,Y etc. and depletion in labile elements like Na, Li etc.
(McArthur et al., 2008) suggest this to be a palaeosol that was
exposed and weathered over a long time period. The calcretised part
is rich in carbonates (McArthur et al., 2008). The cumulative curve
and frequency of grain size distribution of palaeosol sample (Fig. 4a)
shows a weak bimodal population, fine sand to coarse silt, very fine
skewed, very leptokurtic nature. All of these suggest a breakdown of
host rock mineral during the pedogenesis. However, low silt
percentage suggests a rather limited soil formation process possibly
in a low rainfall regime (Ellis, 1980). Due to low TOC content samples
from this facies could not be dated in the present location but these
are certainly younger than 23 ka and older than 9 ka (Fig. 3). Elsewhere in the western delta plain (viz. Kolkata and Diamond
Harbour) this facies has been variously dated as w24 ka and 14 ka
(Fig. 7; Hait et al., 1996a; Stanley and Hait, 2000).
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A. Sarkar et al. / Quaternary Science Reviews 28 (2009) 2564–2581
Fig. 3. Facies types and graphic log correlation along SW1-AP transect at Barasat; different facies associations and stratigraphic units are shown; also shown d13C profiles against
individual log.
Facies D is represented by gray coloured medium to fine grained
sands. The sand:mud ratio is quite high. It erodes deep into marine
facies A resulting in a geometry of low width:depth ratio. From the
subsurface data its contact with laterally adjacent sandy facies B
could not be characterized. Nevertheless comparison of lithologs
and ages between the two facies shows that while facies B is
certainly older than palaeosol (>24 ka), the OSL dates of quartz
grains in facies D provide 23–17 ka (near-LGM) age. Further, the
chemical compositions of both sands are indistinguishable (McArthur et al., 2008). Together these evidences suggest that this facies
also eroded the pre-LGM facies B and re-deposited the sands as
facies D. Much less concentrations in most major and trace
elements compared to younger sands of facies E (McArthur et al.,
2008) suggest intense reworking and leaching similar to facies B as
mentioned earlier. Based on the evidence above we interpret this as
an incised fluvial valley fill deposit during the LGM.
5.4. FA4: Early Holocene estuarine valley fill/aggrading fluvial
channel sand, overbanks and peat swamps
The FA 4 environment consists of facies E and F. Facies E is represented by medium to fine grained gray sand with lower sand:mud
ratio than facies B or D in general. Towards lower part sands are more
amalgamated. But it shows more pronounced fining upward
sequence (than B/D) with less amalgamation of sands up-section
(e.g. core SW 1 and 5; Fig. 3) much alike isolated transgressive
channel fill ribbons embedded in clays (Wright and Marriott, 1993).
This facies is either capped by late Holocene mud (G2) of FA 5 or
laterally interfingers with facies F. A close look at the OSL and 14C
dates across the lower and upper contacts of this facies indicates that
a significant time gap of w10 ka exists at D/E contact while the ages
gradually decrease across the E/F contact upward without any
signature of hiatus (Fig. 3). While the C, D (FA 3) is entirely LGM in age
the E and F FA 4 are essentially Holocene. The spatial disposition,
depth and age data (Fig. 3) indicates that lower part of this facies
occupies the top of the LGM incised valley fill facies D. We infer that
while the lower part of this facies represents estuarine valley fill, the
upper part corresponds to aggrading fluvial channel-overbank
system. The OSL dates (e.g. 7.6–7.1 ka in AP; Fig. 3) suggest that this
aggradation was rapid often depositing >10 m sand in less than a ka.
Its distinct geochemical signatures e.g. high calcite (0.5%), TOC (0.1–
0.5%), and enrichment in most major and trace elements (McArthur
et al., 2008) also suggest fluvial aggradation in a high water table
condition whereby the redox potential was such that most major
and trace elements could not be leached out.
Facies F is represented by gray silt to mud with well developed
peat layers. Towards the southern and northern ends of study area
it overlies the channel sand/estuarine facies E while in the central
part it overlies the LGM palaeosol facies C (Fig. 3). Average TOC
content is highest among all the facies (0.5–1%) but it reaches up to
w34% in peat layers (McArthur et al., 2008). The peat layers with
compressed decayed leaves are found at 17–16 m corresponding to
14
C age of 8 ka. The peat layers are, however, discontinuous in
nature. The base of this facies has an estimated age of w9 ka where
salt tolerant benthic foraminifer like Ammonia is found (also found
at Kolkata, Kolaghat and many other places). Almost at the same or
slightly younger level rich core mangrove pollen (e.g. of Heritiera,
Bruguiera, Acrostichum etc.) are also observed. Temporally the base
of this facies is closer to the estuarine facies E (w9 ka). This, along
with the presence of foraminiferas, suggests a rapid marine incursion during early Holocene time. The grain size analysis of mud
sample just below the peat layer shows a strong bimodality, fine
skewed, platykurtic nature (Fig. 4b) indicating mixing of different
sediment sources. A rapid deposition in a high energy regime and
scouring of older sediments might be responsible for this grain size
distribution. On the other hand grain size analysis of the fine sand/
coarse silt mid-Holocene sample, above the peat layers, show
a unimodal, symmetric, mesokurtic nature consistent with aggradational part of a sequence (Fig. 4c). Together facies E and F suggest
an initial estuarine to later aggrading channel-overbank
A. Sarkar et al. / Quaternary Science Reviews 28 (2009) 2564–2581
environment. The mangrove pollen indicates swamps in a rising
water table condition.
LGM Paleosol
Cumulative mass
retained (%)
a
100
90 Bimodal, very poorly
80 Very fine sandy
70 coarse silt,
60 Very fine skewed
50 Very leptokurtic
40
30
20
10
0
3.0
-1.0
1.0
5.0
sorted
5.5. FA5: Middle to late Holocene channel-floodplains with minor
peat swamps
7.0
9.0
11.0
13.0
15.0
Class weight (%)
7.0
6.0
5.0
4.0
3.0
2.0
1.0
0.0
-4.0
11.0
16.0
Cumulative mass
retained (%)
100
90 Strongly bimodal, very poorly sorted
80 Mud,
70 Very fine skewed
60 Very platykurtic
50
40
30
20
10
0
7.0
9.0
-1.0
1.0
3.0
5.0
11.0
5.0
4.5
4.0
3.5
3.0
2.5
2.0
1.5
1.0
0.5
0.0
-4.0
13.0
15.0
6.0
100
90 Unimodal, poorly sorted
80 Very fine sandy very coarse silt,
70 Symmetrical
60 Mesokurtic
50
40
30
20
10
0
-1.0 0.0 1.0 2.0 3.0 4.0 5.0
11.0
16.0
6.0
7.0
8.0
9.0
Class weight (%)
7.0
6.0
5.0
4.0
3.0
2.0
1.0
0.0
-4.0
1.0
6.0
11.0
16.0
Late Holocene prograding silt
Cumulative mass
retained (%)
d
100
90 Bimodal, poorly sorted
80 Very fine sandy very coarse silt,
70 Coarse skewed
60 Leptokurtic
50
40
30
20
10
0
-1.0 0.0 1.0 2.0 3.0 4.0 5.0
6.0
7.0
8.0
9.0
Class weight (%)
7.0
6.0
5.0
4.0
3.0
2.0
1.0
0.0
-4.0
The FA 5 environment consists of two facies G1 and G2 occurring as topmost capping unit throughout the area. A lower silty part
(G2; 7–6 ka) with decreased concentration of mangrove pollen and
an upper part with thin fine sand layers interfingering with silts.
Towards the upper part (beyond 6 ka) mangrove pollen are
replaced by terrestrial pollen of C4 grasses like Poaceae, Cyperaceae
etc. (Sen and Banerjee, 1990). This might indicate a gradual
lowering of water table after 6 ka time. Brief, highly localized peat
layers are observed at w7–5 m depth corresponding to 14C age of
3–2 ka. Base of G1 (FA 5) silty sub-facies cannot be distinguished
from the top of F facies (of FA 4) excepting significant change in
biota and pronounced change in d13C values (see following
discussion). Grain size analysis of the fine sand/coarse silt sample
from the upper part of this association (facies G1) show a weak
bimodal, leptokurtic but coarse skewed nature (Fig. 4d) suggesting
higher coarse grain supply than the other two Holocene facies
discussed before. These, along with retreat of mangrove and
occupation of more terrestrial plants, possibly suggest progradational channel/floodplain facies during the late Holocene period. An
environment of aggradational to progradational floodplain with
minor channel sands is envisaged for FA 5.
6. Sediment chemistry and carbon isotope stratigraphy
1.0
Mid- Holocene aggrading silt
Cumulative mass
retained (%)
c
6.0
Early Holocene transgressive mud
Class weight (%)
b
1.0
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1.0
6.0
11.0
16.0
Particle diameter ( )
Fig. 4. Cumulative and frequency curves of grain size distribution for Holocene sediments
of western GB delta. (a) LGM palaeosol, (b) Early Holocene transgressive mud, (c) MidHolocene aggradational fine sand/silt, (d) Late Holocene progradational fine sand/silt.
In general the channel sands have much lower TOC compared to
overbank silt or clay. Also, organic matter (and d13C compositions)
within sands may not reflect the in-situ origin. Because pre-Holocene sediments have much higher sand:mud ratio with thicker
channel sands (facies association FA 2, FA 3) compared to Holocenes
(facies association FA 4, FA 5) it was not possible to retrieve
continuous isotopic signature for the lower part of the successions.
Similar problem arose for the estuarine valley fill and aggrading
channel sands of Holocene age. However, these sands have intervening mud layers which were available for isotopic analysis (e.g.
core SW 1, 5). The most continuous isotope stratigraphy could be
reconstructed in the central part of the study area where the LGM
palaeosol and Holocene sediments are thickest without any channel
erosion (e.g. SW 4 and DP). Isotopic and chemical data including total
carbon, nitrogen and C/N etc. are given in supplementary data Table
S3. Fig. 3 shows d13C values of bulk organic matter plotted against
depths and lithologs for all the six cores. Fig. 5 shows the depthvariation of d13C in SW4, DP and SW5 along with several other
climate proxies. For SW4 and DP only a generalized litholog is
provided in Fig. 5 as the depth-wise facies variations are almost same
at these two locations. The d13C variation in SW 4 and DP are
remarkably similar. In SW4 d13C shows continuous upward enrichment from 23.3 to 18.1& (>5&) within the palaeosol. In DP the
enrichment is slightly less w3&. The d13C shows a rapid and large
depletion (w28&) immediately above the palaeosol at the onset of
Holocene sedimentation between 9 ka and 8 ka. The depletion is
9.6& at SW 4 and w8.5& at DP. Thereafter d13C remains constant (at
w27& level) up to w7 ka. From 7 ka to 1 ka the d13C shows
a second phase of enrichment reaching maximum (w18&) near
the core tops. The enrichment is again very large ranging from 9&
(in DP) to w10& (in SW 4). In cores SW 1, 5, 6 and 7 the isotope
profiles are incomplete due to the presence of number of sand layers
below mid-Holocene. Hence the 9–8 ka depletion is un-recorded in
all of them. These cores, however, record only the middle to late
Holocene enrichment in their clay layers. Even there the magnitude
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A. Sarkar et al. / Quaternary Science Reviews 28 (2009) 2564–2581
Fig. 5. d13C, C/N, TOC profiles of core SW4 and DP. Large depletion above the palaeosol suggests transgression and mangrove domination; also note high sedimentation rate in early
Holocene corresponding to stable d13C value (aggradation) and middle to late Holocene progressive d13C enrichment. For comparison d13C profile of sand rich SW5 core is also
shown. Temporal changes in biota (including pollen) at Kolkata (Sen and Banerjee, 1990) and Chilka lake (re-drawn from Khandelwal et al., 2008) indicate control of vegetation on
d13C variation. d13C–C/N data for Chapai-Nawabganj are re-plotted from Meharg et al. (2006).
of enrichment is large (w7–8&) in core SW 1 and 5, while it is
smaller (w2&) in SW 6 and 7. This is because numbers of intervening clay layers are present towards the base of the sections at
SW 1 and 5 which could be isotopically analysed. SW 6 and 7, on
the contrary, preserved only the topmost clayey part of the
Holocene (Fig. 3). Taken together d13C shows a very consistent
variation from LGM to late Holocene throughout the study area.
Such consistent variation and large changes in d13C in all the cores
suggest a specific causative mechanism that perturbed the carbon
budget. The most common cause in these near coastal settings is
the change in sources of organic matter having wide ranges of
d13C compositions (Megens et al., 2002). Because C/N ratio of
organic matter is also a potential tool for discriminating the
sources of organic matter (Meyers, 1994; Andrews et al., 1998),
a comparison of d13C and C/N profile is necessary at this stage.
Fig. 5 also shows the C/N profiles for core SW 4 and DP along with
TOC profile of SW4. The C/N ratio shows continuous enrichment
within palaeosol from w3 to 12 much alike the d13C. The
maximum C/N value of w20 is, however, found towards the base
of Holocene section. C/N gradually decreases between 9 and 7 ka
following which it remains steady between w7 and 1 ka. Changes
in both d13C and C/N, therefore, strongly indicate changing source
of organic matter during the LGM-Holocene period.
6.1. Fingerprinting sources of organic matter by d13C and C/N
Estuarine or deltaic sediments receive both autochthonous (insitu plant community) and allochthonous (transported either by
river or tidal incursion of ocean water) organic matters. Terrestrial
or aquatic plants, algae and marine particulate organic matters
(POM) have large difference in their d13C compositions essentially
arising due to either the differences in photosynthetic mechanism
or sources of carbon used. Terrestrial plants have two major modes
of photosynthesis namely, C3 and C4. In general plants preferentially fix lighter carbon 12C via diffusion during photosynthesis
making the organic matter highly depleted compared to atmospheric CO2 (average d13C w8&; Keeling et al., 1995). d13C for C3
plants (e.g. large land plants including mangroves) ranges from
23& to 30& while the C4 plants (grasses, shrubs) have average
value of w13& (Meyers, 1994). Freshwater aquatic plants,
however, have large d13C range from 50 to 11& (Keeley and
Sandquist, 1992). C/N ratio of C3 (12; Tyson, 1995) and C4 (30;
Meyers, 1994) often overlap. d13C of freshwater algae (26& to
30&) are substantially negative than marine algae (16& to
23&; Meyers, 1994), the later having a small difference with C4
plant values. So is the case for marine POM (represented mostly by
phytoplanktons) having ranges (21& to 18&; Middelburg and
A. Sarkar et al. / Quaternary Science Reviews 28 (2009) 2564–2581
Nieuwenhuize, 1998) closer to C4 plants. Further, large variation
exists either in isotopic composition or C/N ratio within C3 and C4
communities even at generic level. It is, therefore, prudent to use
both d13C and C/N to fingerprint the organic matter source. Fig. 6a
shows the ranges of d13C–C/N ratio of C3 (including mangrove), C4
plants, lacustrine and marine algae and marine POM from BOB. Also
shown values for mangrove plants and selective C4 grasses (e.g.
Poaceae) commonly found in GB delta.
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Before discussing the use of d13C–C/N for source identification it
is necessary to assess the post-depositional diagenetic effect on the
sedimentary organic matter. Early loss of labile organic component
in vascular vegetation do cause change in d13C but it is minor and
the large differences in source organic matters are usually
preserved particularly in younger sediments (Lamb et al., 2006).
The effect is more in C/N ratio. For example, degradation of
terrestrial organic matter causes enrichment in refractory lignin
thereby increasing the C/N ratio (Fogel et al., 1989; Thornton and
McManus, 1994). In spite of this, the general trend and relative
changes in C/N are found to be preserved even in most dynamic
coastal systems (Lamb et al., 2006). Except the two samples with
very high TOC in the peat layer of DP the average concentration (%)
of total nitrogen (TN) and TOC in our cores are 0.03 0.02 and
0.27 0.16 respectively. These values are exactly similar to average
TN and TOC values measured in modern floodplains of Ganges
(Padma river) in Bangladesh (0.03 0.02% and 0.34 0.23%; Dutta
et al., 1999) suggesting very little diagenetic alteration in general.
Excluding the two peat samples with very high TOC, those might
bias the trend, cross plot between total carbon and nitrogen
(Fig. 6b) shows a linear relationship where the regression line
passes very close to the origin indicating that the both are organically derived (Hedges et al., 1986). The slope of this line translates
to an average C/N ratio of w9 for these organic matters which are
close to lacustrine and marine algae or some C4 plant values. Based
on these evidences we infer that both d13C and C/N have not been
diagenetically altered to any large extent and retained their original
source signatures.
Because marine POM (phytoplanktons) can potentially mix into
the deltaic/estuarine system and C/N ratio often fail to discriminate
terrestrial and marine phytoplanktons or algae an independent
tracer is needed for assessing the marine contribution. Although
could not be analysed for the entire downcore, d15N values at select
levels of core SW 4 (Supplementary data Table S3) shows range from
3.5& to þ2.7&. These values are much lower than the average d15N
value of wþ5.2& of the POM obtained in the Bay of Bengal. Time
series measurements of POM d15N, hydrolysable carbohydrates and
amino acids in northern Bay of Bengal suggest that the average
terrestrial fluxes from Ganges–Brahmaputra are indeed characterized by lower (wþ3.7&) d15N values (Gaye-Haake et al., 2005; Unger
et al., 2005, 2006). We, therefore rule out any significant contribution of marine POM into the deltaic sediments throughout the
Holocene. If true, the d13C and C/N variation must then be due to the
changes in local vegetation pool as a function of time.
6.2. Carbon isotope and vegetation
Fig. 6. (a) d13C and C/N ratios of core C3 mangroves (Rhizopora, Avicennia, Bruguiera,
Heritiera, Ceriops), tidal mangrove (Acrostichum), C4 grasses and herbs (Cheno-Amaranthus, Poaceae) and freshwater plants (Typha); also shown are lacustrine and marine
algae and particulate organic matter from Bay of Bengal. Dotted lines show the total
ranges of C3 and C4 plants. C3 and C4 plants have distinct d13C values but overlapping
C/N ratios; note generally higher and lower C/N ratios for mangroves and C4 plants
(e.g. Poacea and Cheno-Amaranthus) respectively commonly found in Holocene sediments. The data compiled from various studies on modern plants, pollen and marine
organic matters (Rodelli et al., 1984; Hemminga et al., 1994; Meyers, 1994; Cifuentesa
et al., 1996; Amundson et al., 1997; Ehleringer et al., 1997; Middelburg et al., 1997;
Peñuelas and Estiarte, 1997; Stribling and Cornwell, 1997; Hornibrook et al., 2000;
Marchand et al., 2005; Muzuka and Shunula, 2006; Descolas-Gros and Scholzel, 2007).
(b) Correlation between total organic carbon and nitrogen from sediments of western
GB delta.
To retrieve the possible link between isotope stratigraphy and
vegetation change, we compared the pollen obtained in the Kolkata
section (Sen and Banerjee, 1990) where the lithostratigraphy and
sedimentation rate are quite similar to those at Barasat. Fig. 5 shows
the isotope profiles (core SW 4, DP and SW 5 with channel sandclay section) and biotic assemblages in Holocene Kolkata section.
Also shown the ages (ka) based on estimated sedimentation rates
(given as m/103 year) against the lithlogs. The LGM palaeosol is
completely barren of any spore/pollen. Such is the case for the
palaeosol throughout the western delta plain (Hait et al., 1996a;
Stanley and Hait, 2000). This is possibly due to the prolonged
exposure of this soil whereby most biotic elements got decomposed. As mentioned such decomposition can alter both d13C and C/
N. Mostly 13C depletion occurs due to selective loss of the isotopically heavy carbohydrate fractions, but occasional enrichment due
to loss of lipid and concentration of cellulosic material is also found
(Spiker and Hatcher, 1984; Meyers et al., 1995). However, this
change can at best be only 1–2&; also the TOC within palaeosol is
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invariant and hence the large enrichment of w 3–5&, must be
reflecting the change in in-situ organic matter in the palaeosol. The
organic matter in the lower part of the palaeosol, immediately
above the pre-LGM FA3, possibly has substantial C3 component.
Note that elsewhere, like in Kolaghat, the clay from FA3 yielded
mangrove (C3) pollen at w31 ka (see previous discussion). We
attribute the enrichment in the top of palaeosol to the prolific C4
vegetation (mostly grasses) on the western delta floodplain during
the LGM. Indeed both GIS-based vegetation map (Ray and Adams,
2001) and CARAIB dynamic vegetation model show expansion of
tropical grassland (in place of modern tropical seasonal forests) in
the western part of the GB delta often reaching a semi-desert
condition during the LGM (Galy et al., 2008a). Corresponding predicted model based change in average d13C of organic matter
(biomass, litter, soil carbon together) is from 26& (at present) to
w20& during the LGM. The highly enriched (w18 to 20&)
value in the palaeosol facies at Barasat possibly gives a direct
evidence of C4 grassland expansion during the LGM. The correlated
change between d13C and C/N ratio in palaeosol is, however,
intriguing. In general C/N is not a very effective criteria to
discriminate C3 and C4 (Fig. 6a) although some C4 grasses like
Poaceae can have very low C/N (w5–10; Descolas-Gros and Scholzel, 2007) similar to values found in the lower part of the palaeosol.
The most plausible reason of C/N enrichment to the topmost part of
the palaeosol is the degradation and consequent lignin concentration in the organic matter, as mentioned above, due to extreme
weathering atop the palaeosol in an arid climate.
The large and rapid d13C fall at the onset of Holocene (w9 ka),
immediately above the palaeosol, occurs in a zone rich in benthic
foraminifera Ammonia and prolific pollen of arboreal mangrove
plants such as Heritiera, Bruguiera, Acrostichum etc. slightly above it.
Clearly the range of d13C–C/N values of these plants (Fig. 6a) suggest
that the 13C depletion of the order of 8–9& and highest C/N ratio
(18–20) at this level were caused by a rapid increase in mangrove
vegetation. Although presence of foraminifera at this level indicates
a marine input, the total carbon budget was overwhelmingly
controlled by particulate organic matter derived from mangrove
forests that was far more extensive than today. TOC is also high at
this level (Fig. 5). The contact between palaeosol and earliest
Holocene sediment, therefore, demarcates a major change in
vegetational history of the western delta plain. Between 8 and 7 ka
the d13C remains stable at w27& where mangrove vegetation,
represented by Avicennia, Bruguiera, Rhizopora, Ceriops etc.
continued to dominate. The C/N ratio during this time gradually
returns to w11. Using the sediment d13C, a simple mass balance
calculation (Thornton and McManus, 1994) is applied to estimate
the relative contribution of C3 and C4 vegetation in the LGMHolocene section. The fractional contribution of C3 organic matter
is calculated by
period of constant d13C represents an eco-system transition between
pure mangrove to mixed mangrove-C4 vegetation. However, the
pollen of Poaceae starts appearing only after 7 ka with concomitant
decrease in mangrove (C3:C4w50:50). The 9–10& gradual enrichment in d13C and large decrease in mangrove pollen, appearance and
expansion of Poacea, Cyperacea and Cheno-Amaranthaceae between
7 ka and 2 ka in general suggest lowered water table and disappearance of mangrove eco-system. The decrease in C/N ratio attests
the stabilization of C4 vegetation during this phase. Mangrove
pollen briefly appear between 4 and 3 ka (albeit with a much lesser
magnitude) but its effect is not seen on the d13C composition. This
level coincides with a minor peat forming event (high TOC in Fig. 5)
in the western delta plain possibly representing local swampy
conditions within the late Holocene floodplains (Figs. 3 and 5).
Towards the top of the sections at w1 ka the mangrove completely
disappears and C4 grasses like Poacea, Cyperacea reaches their
maxima. The latest Holocene floodplains of the western part of the
GB delta thus show invasion of C4 vegetation (C3:C4w30:70) when
the mangrove front was pushed back to its present location. The
estimation of C3:C4 for the LGM (40:60) closely agrees with those
obtained from both bulk and n-alkane d13C data from Bengal fan
(Galy et al., 2008b) but shows considerable departure during the
early and late Holocene period. In particular the late Holocene
continental sediments show progressively large invasion of C4
community (C3:C4w30:70) not recorded in the marine sediments.
The LGM-late Holocene change in vegetation, observed in pollen
and d13C at Barasat and Kolkata, seems to have affected the entire
Indian coast. A well dated sediment core from the Chilka lake (a
lagoon-barrier bar complex in east coast; Figs. 1a and 5) shows
exactly the similar temporal changes in vegetation (Khandelwal
et al., 2008). Like Kolkata, pre-9.5 ka level at Chilka also shows poorly
preserved pollen of C4 plants like Cyperaceae–Chenopodiaceae and
ferns–freshwater taxa assemblage. Core mangrove assemblage like
Rhizophoraceae–Avicennia–Sonneratia–Excoecaria rapidly expands
at 9.5 ka and continues up to 7.5 ka. From 7.5 ka to 2 ka mangroves
decreased, remained at a low abundance and were replaced by C4
assemblage like Poaceae–Cyperaceae–Chenopodiaceae and hinterland taxa. Between 2 ka to present mangroves completely disappeared and C4 plants had taken over. On a larger spatial scale the
pattern of d13C and C/N change from Barasat has striking similarity
with those from Chapai-Nawabganj (Fig. 5; data re-plotted from
Meharg et al., 2006), a northern location in Bangladesh although the
sampling resolution is poor. Sedimentation rate at this location is
very high and 14C date at w25 m depth is w5 ka (BGS and DPHE,
2001) yet the d13C change reflecting C3–C4 transition, as observed in
Barasat during the Holocene, is also present here. Such uniformity in
d13C stratigraphy suggests that the vegetation change was indeed
ubiquitous throughout the western part of the GB delta.
7. Sequence stratigraphy
13
13
13
d Cbulk ¼ f C3 d CC3 þ f C4 d CC4
where, d13Cbulk ¼ sample value, fC3 ¼ C3 fraction, fC4 ¼ C4 fraction,
d13CC3 (28&) and d13CC4 (13&) are average values for C3 and C4
plants respectively (Fig. 6a; Meyers, 1994). Estimation shows that
while the C4 dominated the LGM section (C3:C4w40:60), earliest
Holocene transgressive phase was exclusively C3 dominated
(C3:C4w90:10) which was responsible for extensive peat deposit
throughout the region. The 8–7 ka period of near-constant d13C and
slightly decreased fC3 (C3:C4w80:20) possibly indicates no major
change in water table and a gradual return to a mixed C3–C4 vegetation. Because w80% of all C4 species are Poaceae grasses and dicots
like Chenopodiaceae, Amaranthaceae etc. (Keeley and Rundel, 2003)
whose C/N ratio are rather low (<10; Peñuelas and Estiarte, 1997;
Keeley and Rundel, 2003; Descolas-Gros and Scholzel, 2007), this
Marine influence in a low gradient fluvio-deltaic setting can be
traced up to w200 km inland (Posamentier, 2001). Hence, the key
to understand regional sedimentology and build a sequence
stratigraphic framework in a distal fluvial- shallow marine interactive setting lies in the recognition of regionally correlative
surfaces, various systems tracts and their physical position within
the sequence (Bhattacharya, 2006). To achieve this goal attempts
were made to visualize the depositional architecture of the western
delta plain from pre-LGM to late Holocene period through the
correlation of eight region-wide representative sections (Fig. 7).
The sections are spread in a near N–S transect from Dankuni in
north to Bakkhali on the present coastline (see Fig. 1a). While two
sections are from present study area, rests are taken from published
literature. Fundamental to the sequence analysis, however, is the
A. Sarkar et al. / Quaternary Science Reviews 28 (2009) 2564–2581
Fig. 7. Regional graphic log correlation across the Pre-LGM-Holocene sections from western delta plain; log locations are given in Fig. 1a, b. Data for AP and SW4 are from present study; the distance between AP and SW4 is not to the
scale and exaggerated for clarity. Data source for other locations: Sen and Banerjee (1990); Hait et al. (1996a, b); Stanley and Hait (2000). Correlative surfaces viz. unconformity, transgressive surface of erosion and maximum flooding
surfaces are marked. Note thin early Holocene Foraminifera-peat layer rich muds above the FSST (palaeosol-incised valley fill) of LGM age indicating an abortive TST due to very high sediment supply.
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A. Sarkar et al. / Quaternary Science Reviews 28 (2009) 2564–2581
Fig. 8. (a) Facies associations and sea level in the GB delta. (b) Sea-level curve around Indian coast; data source: Banerjee (1993); Vaz (1996); Vaz and Banerjee (1997); Vaz (1999);
Farooqui and Vaz (2000); Vaz, 2000; Pandarinath et al. (2001); Rao et al. (2003); Mathur et al. (2004); Hameed et al. (2006). Note rapid rise between 14 ka and 7 ka; rate of rise
decelerates after 7 ka; sea level was higher-than-modern between 4.5 ka and 2 ka. Red sea and global sea-level curves from Siddall et al. (2003) and Peltier (2002) respectively. pCO2
values between LGM and present are from Smith et al. (1999). (c) Monsoon proxies (G. bulloides upwelling index from Arabian sea, Gupta et al., 2003 and d18O precipitation index in
stalagmite from southern Oman, Fleitmann et al., 2003) over the last 14 ka. Note monsoon intensification between 10 ka and 7 ka; monsoon was weaker at 11–14 ka and late
Holocene (post-6 ka). Denudation rates in higher Himalayas was simultaneously higher (w5–6 times) than late Holocene (Bookhagen et al., 2005). BOB d13C data from Fontugne and
Duplessy (1986).
facies analysis, identification of discontinuity surfaces and isotope
stratigraphy obtained from the eight cores in the present study area
(Fig. 3). It is important to mention that the subsurface sections both
at Barasat and along the Dankuni–Bakkhali transect may not be
strike perpendicular which often limits our spatial correlation of
facies. Nevertheless it was possible to identify the major depositional changes and surfaces.
Two surfaces served the purpose as entire facies motif could be
constructed considering them as reference: (i) the conspicuous
lithological contact coinciding with the boundary of pre-LGM sand
or clay of FA 2 and its overlying palaeosol (facies C). At Kolkata the
palaeosol yielded a date of w24 ka (supp. Table S1) that might
represent an integrated age over a large period of soil development.
Tracing laterally, e.g. at Barasat, this contact is placed at the
boundary between the marine clay (facies A) and 23 ka sand of
facies D. (ii) the surface atop the palaeosol demarcated either by
few meter thick benthic foraminifera rich clay layer or peat layer
rich in mangrove pollen.
Regionally, occurrence of TOC rich marine FA 1 at the base of all
the sections suggests its deposition like a marine onlap in a transgressive mode (Galloway, 1989). McArthur et al. (2008) reports
a minimum 14C age of w27.2 ka for this facies. The shallowest depth
at which the upper contact of this facies found is w45 m below
present msl (Fig. 3). Hence the sea level, during which this transgressive clay deposited, must have been several tens of meters
higher than this depth. Because the tectonics induced change in
relative sea level in western delta plain is minor (Goodbred and
Kuehl, 2000a, b) and very high sea level only existed during the last
A. Sarkar et al. / Quaternary Science Reviews 28 (2009) 2564–2581
interglacial stage (Siddall et al., 2003), we infer that this clay unit
was deposited in marine oxygen isotope (MIS) stage 5 (pre-80 ka)
only. The exposure of this unit was possibly caused by relative sealevel lowering between MIS 5 and 4 when sea level was 90–100 m
lower than today. The absolute sea-level fall between MIS 5 and 4
was of course w70 m (Siddall et al., 2003). Fig. 8a shows the
possible timing and duration of FA1 against the sea-level curve of
Siddall et al. (2003).
Up the stratigraphic section the fluvial channel sand-overbank
deposit (FA2) overlies the shelfal FA1. The nature of contact or
transition between FA2 and FA1 could not be studied due to lack of
data (present only at DP or AP; Figs. 3 and 7). However, between
Barasat and Kolaghat the oxidized sand atop the FA1 laterally
interfingers with w31 ka clay layers (Fig. 7) yielding mangrove
pollen possibly indicating local swamps developed during a later
highstand. The spatial variation in sand:mud ratio is noticeable in
this unit e.g. high at Barasat but low at Kolaghat. Because early
highstand fluvial channel sands are essentially isolated and
engulfed in mud while the late ones are more amalgamated with
high sand:mud ratio (Shanley and McCabe, 1993) the regional
variation of this fluvial facies suggests preservation of different
components of highstand sediments. Together the age, biota and
facies relationship suggest that deposition of this highstand fluvial
sand-clay unit took place during the MIS stage 3 (50–30 ka) when
sea level was higher than MIS 4 but w50–60 m lower below
present (Siddall et al., 2003; Fig. 8a).
There exist some data on sea-level change around Indian coast
between MIS 2 and late Holocene which need to be assessed before
discussing LGM-Holocene sequences. Fig. 8b shows the chronologically well constrained sea-level data plotted against age
between 20 ka and recent. While only few data are available on the
basis of dated gastropod or oyster shells in the west coast, east coast
is relatively better constrained in terms of data frequency (well
dated corals or bivalves). A cubic spline, fitted through the data
points, shows that the sea level was indeed lower by >100 m
during the LGM (20–18 ka). For comparison oxygen isotope based
sea-level curve from Red sea (an area closer to the GB delta; Siddall
et al., 2003) and theoretically estimated global sea-level curve
(Peltier, 2002) are also plotted in Fig. 8b. The global sea-level curve
closely matches the coral based curve of Barbados (Fairbanks,
1989). Although general pattern of Indian sea-level curve agrees
with Red sea or global curve, there is also considerable mismatch
between them. This could partly be due to weak Indian data base.
Between 18 and 14 ka the sea level was low with very slow rate of
increase corresponding to the late lowstand. Subsequently the sea
level starts rising rapidly from aroundw14 ka, the fastest rise
occurring between 14 and 7 ka (w1.5 cm/year). Rate of sea level
change decreased after 7 ka (w0.7 cm/year), maximum sea level
(w5 m higher than present msl) reaching at 4.5 ka. A sea level
standstill occurs between 4.5 ka and 2 ka after which it comes
down to the modern level. The depositional sequence of GB delta
between LGM and Holocene can, therefore, be viewed against these
eustatic changes.
Development of meter thick palaeosol with calcretes (facies C)
atop the highstand FA 2 suggests maximum subaerial exposure in
the interfluve areas. The low TOC and d13C values suggest presence
of grassland vegetation and intense oxidation in a possible dry
environment. The upward trend of d13C within the palaeosol (Fig. 5)
suggests that vegetation gradually transformed from a mixed C3–
C4 in the lower part to dominant C4 grassland towards the top. This
along with available ages (24 and 14 ka at Kolkata and Diamond
Harbour respectively) suggests that possibly the soil formed at the
culmination of extreme progradation of fluvial sands of MIS 3
highstand and continued till 14 ka. The age bracket of the palaeosol
corresponds to the MIS 2. We postulate that the palaeosol was
2575
formed during the maximum sea-level fall of >100 m during the
LGM (Fig. 8b). Important to note that the calcretised part of this
facies is quite thinner indicating either substantial erosion of soil
itself during the base level fall or not so extreme aridity (Tandon
and Gibling, 1997; Kraus, 1999) during the LGM. Grain size analysis
too indicates the soil to be rather immature. Considering the
average age of palaeosol in western delta plain as w20 ka and the
oldest age of the overlying sediment (facies F) is w9 ka the contact
between FA 3 and FA 4 signifies a hiatus of w104 years. The
geometry of the palaeosol suggest that they occupy top of buried
uplands or interfluves (Figs. 3 and 7). Laterally the palaeosol is
deeply incised. Such deep fluvial incision is characteristic of large
sea-level fall of >110 m as documented in Java shelf (Posamentier,
2001) and possibly occurred at the nadir of the sea-level fall of
>100 m during the LGM or MIS 2. Sequence stratigraphic model
suggests that such deep incision as much as 40–70 m above the
shelf edge preferably occurs when the river entrenches a channel
through a subaerially exposed convex-up topography like highstand prism (Talling, 1998). This is analogous to the present study
area where the incision takes place within the MIS 3 highstand.
Together the palaeosol and incised valley make an ancient buried
landscape forming upland terraces and valleys. The correlative
contact (thick dashed line in Fig. 7) between MIS 3 sand and
palaeosol on the uplands or FA1 and facies D (base of incised valley
fills) is of regional importance and implies a major unconformity
laterally traceable in the entire delta plain. The LGM incised valley
was filled by sands of 23–17 ka (Facies D) age possibly during the
early to late lowstand when the rate of sea-level fall decreased.
The palaeosol is immediately overlain either by benthic foraminifera rich clay-silt (facies F) horizon indicating marine incursion
(possibly during rapid post-LGM sea-level rise; Fig. 8b), e.g. at
Diamond Harbour or isolated clay-embedded transgressive channel
fill/estuarine ribbon sands (facies E) in incised valley (e.g. at Barasat), both having similar early Holocene age. The age difference
between incised valley fill (23–17 ka) and estuarine valley fill
(w8 ka) suggests substantial scouring at the contact in between,
either by estuarine channels or marine ravinement surfaces (Allen
and Posamentier, 1993; Rossetti, 1998). Such transgressive reworking also left its traces in the grain size distribution showing a strong
bimodality where coarser population is mixed with a finer one
possibly derived from two different sources. The spatial disposition,
depth and age data (Fig. 3) indicate that lower part of the facies E
sands occupy the top of the LGM incised valley fill facies D (Figs. 3 and
7). Such occupation of partly filled incised valley fills suggests that its
upper part has possibly been converted as an estuarine valley fill at
the onset of transgression (Dalrymple et al., 1992). The foraminiferal
layer above the palaeosol is thin and interpreted as ‘‘event concentration’’ during rapid sea-level rise (Cattaneo and Steel, 2003).
Nevertheless an apparent diachroneity (Cattaneo and Steel, 2003) of
this layer is identifiable. For example, at a distal location like Diamon
Harbour the age of the base of this layer (assuming a sedimentation
rate of w4.4 m/ka; Fig. 2) is 9 ka while at the most inland location at
Dankuni it is w8 ka. In absence of any identifiable back-stepping
stacking pattern in subsurface cores, this perhaps suggests typical
retrogradational nature of the transgression.
Towards inland the foraminiferal horizon nearly merges with
the peat layers. d13C and pollen at this level indicates complete
occupation by C3 mangrove vegetation eventually giving rise to
peat deposits. Because peat formation takes place when the water
table is at its highest (accommodation > sediment supply) and
marine maximum flooding surface (MFS) overlaps with the peat
layers (Hamilton and Tadros, 1994) we consider the top of the
earliest Holocene w9–8 ka old foraminifer-peat horizons as the
MFS in western delta plain (thick dotted line in Fig. 7). Likewise we
assign the base of this unit (either between facies C and F or D and
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A. Sarkar et al. / Quaternary Science Reviews 28 (2009) 2564–2581
E; Fig. 3) as the transgressive surface of erosion (TSE; thick solid
line, Fig. 7; Posamentier and Vail, 1988). Bounded between the two
key surfaces, the LGM unconformity below and early Holocene TSE
at the top, the palaeosol-incised valley of LGM are interpreted as
a forced regressive unit, a product of Falling Stage Systems Tract
(FSST) following revised Exxon systematic (Hunt and Tucker, 1992;
Catuneanu, 2002).
Association of forced regressive unit like development of
palaeosol and coeval channel incision allowed us to interpret the
relative sea-level fall a result of forced regression (Hunt and Gawthorpe, 2000) and assign type 1 character for the unconformity
(Catuneanu, 2002). Between TSE and MFS lies the transgressive
systems tract (TST; 9–8 ka) either forming retrogradational estuary
in the upper part of the incised valley fill or peat swamps in the
adjacent floodplains. Important to note that on a regional scale
(Fig. 7) the thickness of the TST (in interfluves) is very low (1–2 m),
does not vary much and the typical proximally thinning wedge
shape geometry is barely discernible (Fig. 7). The change in sedimentological character across the MFS is minor from dominant
mud in the TST to fine sand to coarse silt above it. But near-constant
d13C value with stabilized mangrove vegetation between 8 and 7 ka
or rapid estuarine valley fillings (Fig. 5) suggest an aggradational
stage in a constant water table condition. The unimodal, mesokurtic size distribution (Fig. 4c) also indicates insignificant change
in energy level. The aggradational thickness is also much higher (4–
5 m) than the TST. This is in spite of the fact that the sea-level rise
was still rapid during this time period (Fig. 8b). Because effects of
tectonism are negligible in western delta plain such thick aggradation during an early transgressive event can be attributed to
a very high sediment supply (see later discussion). This possibly
hindered the development of a full fledged thicker retrogradational
stacking pattern and instead produced an aborted TST. We, therefore, place the TST-HST contact at the base of this aggradational
sequence unlike Goodbred and Kuehl (2000a) who interpreted it as
‘‘TST aggradation’’.
Further up in the stratigraphic section fine sand to coarse silt
deposition continues between 7 ka and 1 ka. The amount of
channel sands in this segment are only minor, engulfed in mud/silt
as encountered at Haldia and Bakkhali region (Fig. 7). The d13C
indicates a gradual decrease in C3:C4 ratio possibly as a result of
increased progradation, lowering of relative sea level and pushing
of the mangrove front towards the present day coastline. The
aggrdational part just above the TST (8–7 ka; Fig. 5) and the progradational part (7–1 ka) above it possibly represent the early and
late highstand respectively representing the mid- to late Holocene
highstand systems tract (HST). The supply of coarser fraction
increased during this time with the formation of minor channels.
Alternatively, the less frequency of the channel sands at the top of
HST could possibly be due to their low preservation potential
during the ensuing sea-level fall (Goodbred, 2003). Localized peat
layers re-appear between 3 and 2 ka. This coincides with a sea level
higher than the present msl (Fig. 8b). However, unlike earliest
Holocene it is not identifiable in d13C signal; nor they occur as
regionally persistent horizon. Nevertheless a brief period of high
water table can be envisaged for this period.
8. Evolution of western delta plain
8.1. Last interglacial (MIS 5) to interstade (MIS 3): 125–24 ka
Extensive deposition of marine clay (FA1) took place during the
MIS 5 highstand. Similar w125 ka old marine clay has also been
observed in the Mahi basin, western India (Juyal et al., 2000)
suggesting it to be a regional phenomenon. Following the MIS 5
highstand and subsequent lowstand of MIS 4, fluvial channel
floodplain deposition took place during the MIS 3 highstand.
Typical mangrove forests in swampy or estuarine system grew
under a high monsoon rainfall condition as recorded by various
climate proxies. Earlier studies indicated large glacier advancement in higher Himalayas during MIS 3 (Owen et al., 2002).
Because glacier cover decreases the land-sea temperature gradient
via albedo, it posed an apparent contradiction with high monsoon
regime (Goodbred, 2003) during this time. Recent 10Be and 21Ne
chronologies of erratic boulders from moraines of monsoonal
Himalayas in southern Tibet suggest that all the MIS 3 ages
actually represent the ages of exhumation-denudation in a high
rainfall regime and not the glacial advance (Schaefer et al., 2008).
This is consistent with the records of stronger upwelling in the
Arabian sea (Prell and Kutzbach, 1987). Elsewhere in India
extensive fluvial (both channel and floodplain) aggradation under
a high rainfall regime has been observed during MIS 3 (50–25 ka),
viz. lower Narmada valley (Bhandari et al., 2005), Maharashtra
(Kale and Rajaguru, 1987), southern Gangetic plains (Gibling et al.,
2005; Williams et al., 2006; Sinha et al., 2007) and margins of Thar
desert (Juyal et al., 2006). Depositional and climatic data for MIS 3
are absent in the plains of GB delta but high carbonate content
from the submarine fan core (Weber et al., 2003) in BOB was
earlier interpreted as reduced terrigenous flux in MIS 3. Numbers
of evidences suggest that carbonate content during MIS 3 was high
in northern Indian ocean mainly due to higher biogenic productivity (Sarkar et al., 2000). Hence the high carbonate in MIS 3 fan
sediments too possibly reflect increased productivity and not
reduced siliclastic flux. Actually the oceanic productivity increase
might have masked the enhanced terrigenous flux from the river.
The HST channel sand and mangrove pollen rich clays (FA 2) testify
such sediment trapping in a high sea-level warm strong
monsoonal environment.
8.2. Last Glacial Maximum and lowstand (MIS 2): 24–14 ka
MIS 3 was followed by sea-level lowering of >100 m, extensive
palaeosol (FA3) development on the exposed floodplains of GB delta
and sediment bypassing to the Bengal during MIS 2. This palaeosol
has been encountered throughout the GB delta (see earlier discussion). In western India 24–26 ka old palaeosols with pedogenic
calcretes have also been observed in Narmada (Allchin et al., 1978)
and Mahi (Maurya et al., 2000) river basins. The sea-level minimum
in LGM caused the maximum valley incision marking the end of FSST.
OSL dates of w20 ka at the base of a valley fill deposit in coastal
Narmada basin (Juyal et al., 2006) suggest that such incision was
ubiquitous around the entire Indian coast. In GB delta the OSL dates
indicate that at least a part of the incised valley was filled up by sands
(of FA3) during the lowstand between 18 and 14 ka.
Numerous studies have indicated that summer monsoon was
simultaneously weak in MIS 2 with lesser discharge through GB
system and higher salinity in BOB (Cullen, 1981; Duplessy, 1982;
Sarkar et al., 1990). Fig. 8c shows climate interpretation over the last
14 ka based on various proxies. Trends of two monsoon proxy
records viz. Globigerina bulloides upwelling intensity from Arabian
sea (Gupta et al., 2003) and d18O precipitation record in stalagmite
from southern Oman (Fleitmann et al., 2003) suggest weaker Asian
monsoon in pre-Holocene period. The enriched d13C data in Bengal
fan during the LGM was earlier interpreted as lesser input of
terrestrial organic matter into the marine system due to weaker
monsoon and Ganges discharge (Fontugne and Duplessy, 1986).
However, the LGM enrichment in organic matter d13C must have
a strong component of C4 vegetation that thrived onland as shown
in the present work as well as d13C records from Bengal fan (Galy
et al., 2008b).
A. Sarkar et al. / Quaternary Science Reviews 28 (2009) 2564–2581
8.3. Post-LGM transgression and monsoon intensification (MIS 1):
(14–7 ka)
Rapid post-LGM sea-level rise between 14 and 7 ka (Fig. 8b)
pushed the coastline and mangrove front about 100 km inland
(present work; Allison et al., 2003). The lowstand surfaces (palaeosol-incised valley fill) were flooded forming a TST deposit between
9 and 8 ka. On interfluve areas the TST deposited a foraminifera rich
silt/mud at its base (TSE) or peat layers decomposed from mangrove
plants above (MFS). Incised valleys were covered by estuarine sands.
The estimated sediment accumulation rate during this time was 4–8
times higher (see discussion earlier) than present (Fig. 5). Such large
increase in early Holocene sedimentation rate has also been
observed in the eastern part of GB delta (Goodbred and Kuehl,
2000a, b; Goodbred, 2003) and explained by intensification of
monsoon and consequently much higher water/sediment discharge
through GB river systems. The climate proxy records show highest
monsoon intensity between 10 and 7 ka following which its strength
gradually decreased (Fig. 8c). The organic matter d13C in BOB nearly
follows the trends of monsoon variation. Because the erosion rate in
the Himalayan hinterland is directly correlated to intensity of rainfall
(Singh et al., 2008) and satellite based modern rainfall measurements show orographic rainfall maximum along the stretch of
southern Himalayan margin (Bookhagen and Burbank, 2006), high
monsoon intensity during the Holocene must have an imprint of
erosion in sediment record. Indeed measurement of sediment
accumulation rate in higher Himalayan lakes associated with large
landslides show much higher denudation rate (w4.3 mm/year)
between 10 and 7 ka that was driven by high precipitation, pore
water pressure, slope failure and eventual mass wasting. The late
Holocene to modern denudation rate is w0.7 mm/year, about w5–6
times lower (Fig. 8c; Bookhagen et al., 2005). Interesting to note that
magnitude of changes in Himalayan mass wasting rate are quite
close to the change in sediment accumulation rates between early
and late Holocene measured in the western delta plain (see Fig. 2).
This strongly suggests that large early Holocene increase in sedimentation rate was essentially controlled by monsoon induced
increase in sediment flux in the hinterland.
We postulate that the transgressive invasion possibly took place
along the paleo-valley (palaeo-Ganga?) presently occupied by the
Fig. 9. Frequency distribution of Holocene peat deposits around Indian coast; note
concentration at two time intervals viz. 8–7 ka and 5–4 ka. Data source: Umitsu (1993);
Islam and Tooley (1999); Goodbred and Kuehl (2000a, b); Allison et al. (2003) (East
coast and GB delta), Farooqui and Vaz (2000); Pandarinath et al. (2001); Narayana
(2007) (West coast).
2577
Hoogli River (Fig. 1a). Because the monsoon driven sediment flux at
the hinterland increased at least by 5 times, the newly created
accommodation space was rapidly filled up making the retrogradational TST thickness small (w1 m) but producing a thicker
(w4–5 m) aggradational sequence atop. In the eastern part of GB
delta a much thicker (w50 m) aggradational sequence has been
reported by Goodbred and Kuehl (2000a, b). The GB delta, therefore, considerably differs from other deltas like Mississippi both in
depositional stacking patterns and climate driving force dominating over sea-level rise. This has been clearly demonstrated by
non-linear numerical models of Goodbred et al. (2003) which
shows that the deltas, related to large tropical rivers, responds in
a much stronger way to mega-climate system like summer
monsoon than sea level alone. In fact the climate can drive the
sequence development along the entire fluvial system from its
source to sink as documented in both upper Ganga (Gibling et al.,
2005) and lower delta plains.
The sequence stratigraphic rationale presented above does not
corroborate the earlier inference (Goodbred and Kuehl, 2000b) of
flooding of the LGM lowstand surface as early as 11–10 ka throughout
the GB delta. The evidence of 11–10 ka transgression comes from
dated wood and plant remains from a sediment core in Khulna
(Bangladesh) having the same latitude as the present study area
(Umitsu,1993). TST date older than 9 ka has, however, not been found
in the entire western delta and as far south as Chilka lake in the
eastern coast of India. It is, nevertheless, possible that the accommodation space in the eastern part of the delta had a component, in
addition to sea level, of tectonic subsidence (Goodbred and Kuehl,
2000b; Goodbred et al., 2003) which caused an earlier transgression.
On the other hand our inference of aborted thin TST due to monsoon
driven sediment discharge and rapid aggradation as early as 8 ka is in
agreement with those of Goodbred et al. (2003) indicating it as
a regional phenomenon. In any case the initiation of GB delta as
a whole took place at least 1 ka earlier than the mean formation age of
w7.8 ka of world’s major deltas (Stanley and Warne, 1994).
8.4. Mid-late Holocene progradation and weaker monsoon (MIS 1):
(7–1 ka)
This period coincides with reduction in monsoon intensity
and erosion in higher Himalayas (Fig. 8c). Rate of sea-level
change simultaneously decreased after 7 ka. Sedimentation in
delta plain experienced a progradational phase. This indicates
that rate of sediment supply was higher than the rate of development of accommodation space in spite of the lesser sediment
supply from hinterland. Allison et al. (2003) demonstrated that
the progradation of lower delta plain occurred in five phases. The
earliest progradation occurred at w5 ka in Indian part. As the
accommodation was filled up the major delta lobe formation
switched from west to east culminating at w0.2 ka at the mouth
of the present active Ganges in Bangladesh. Our data from the
western delta plain suggest that following the transgressive shift
of coastline and mangrove front, delta progradation in the
western delta plain took place at least from 7 ka onwards
contrary to what suggested by Allison et al. (2003). The paleoGanges or river Hoogli was the main sediment supply conduit
then. As the progradation was complete Hoogli was abandoned
and the supply occurred through more easterly channels like
Ganges and Brahmaputra. Such shift was attributed to tectonics
in the eastern part of delta (Goodbred, 2003; Allison et al., 2003)
thus making the GB delta considerably different compared to
other major deltas like Mississippi or Nile where the sediment
supply conduit did not significantly change.
The higher-than-modern sea level between 4.5 ka and 2 ka in
a weak monsoon-low sediment supply condition might have
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A. Sarkar et al. / Quaternary Science Reviews 28 (2009) 2564–2581
caused a slight increase in accommodation space. In western delta
plain this might have been the reason of temporary return of
mangrove front between 3–2 ka (e.g. at Kolkata) or local peats (e.g.
at Barasat, see Figs. 3 and 5). Fig. 9 shows the frequency of
mangrove peat layers plotted against age around the Indian coast. It
shows that earliest mangroves appeared in the west coast well
before 10 ka. The east coast mangroves show two distinct peaks one
at w8–7 ka and the other 5–4 ka. The 8–7 ka peats clearly correspond to the early Holocene transgression discussed above.
Although the distribution might have substantial sampling bias,
distinct possibility exists about a later phase of mangroves when,
between 4.5 ka and 2 ka, the sea level was w5–10 m higher than
present. However, the resulting accommodation space must not
have been very large since the mangrove swamps formed only
locally and also their relative abundance remained subdued as
indicated by both pollen and carbon isotopes (Fig. 5). In general sea
level, and not climate, was the major driving force for sequence
development in the penultimate 7–1 ka period, thus largely
producing a progradational HST (FA5) in the western delta plain.
9. Driver of vegetation change
The simultaneous change in local climatic (monsoon) condition, depositional environment and sequence development had
profound impact on the terrestrial vegetation, namely the
proportion of C3–C4 vegetation on the delta plain. This is contrary
to the observation made by Galy et al. (2008b) who, based on
ocean core data from Bengal fan, suggested that the terrestrial
vegetation in Himalayan basin or on Gangetic plains unidirectionally changed from essentially C4 during the LGM to a mixed
C3:C4 up to 8 ka mainly due to the change in atmospheric pCO2
and climate. Thereafter mixed C3–C4 vegetation continued till late
Holocene. Although the origin of isotopically enriched C4 plants
(efficient CO2 users compared to C3 plants) has long been associated with a general decline in atmospheric CO2 during the Cenozoic (Cerling et al., 1997), recent studies suggest that natural
selection of C4 plants is favored by increased perturbation of the
ecological niches (namely fire in an arid climate; Keeley and
Rundel, 2005 or warm season precipitation like monsoon; Bond
et al., 2005) rather than pCO2 change (Huang et al., 2001). Further,
C4 plants are better adapted to a dry water-stressed condition
than C3 plants (Bond et al., 2005). Our work suggests that vegetation was controlled by specific ecological niches created largely
by changes in depositional environment and climate. The dominantly C4–C3 transition was never unidirectional on delta plain,
rather fluctuated at least twice, once from LGM C4 to early
Holocene C3 mangrove (C3:C4 changing from 40:60 to 90:10) and
again from mid-Holocene mixed C3–C4 to late Holocene C4 (C3:C4
changing from 50:50 to 30:70). While the estimated C3:C4 ratio of
LGM at Barasat (40:60) is similar to that of Galy et al. (2008b) viz.
45:55, the late Holocene ratios are quite different (30:70 at Barasat
against 75:25 estimated from Bengal fan). Further, unlike at Barasat, the Bengal fan data suggest that the ratio remained constant
for the last 8 ka (Galy et al., 2008b). The late Holocene C4 invasion
found in GB delta plain or coastal lagoons like Chilka are not
visible in ocean core. The variation in C4 between LGM and
Holocene, thus, cannot be explained by atmospheric pCO2 as the
global pCO2 singularly increased from w190 ppmv to >270 ppmv
during this time (Smith et al., 1999; Fig. 8b). The LGM or late
Holocene increase in C4 is better explained by a combined effect of
weaker monsoon intensity and low water table condition (waterstress) due to relative sea-level lowering. It looks that the ocean
core data provide only an integrated picture not recording the
specific history of trapping and dispersal of terrestrial organic
matters formed by different photosynthetic mechanisms. Indeed
the vegetation model estimates of Galy et al. (2008b) indicate
considerable mismatch (with observed one) in the magnitude of
d13C change between LGM and Holocene. A detail compound
specific isotopic fingerprinting of different biomes and their fate
into soil and sediments in the upper Gangetic plain and GB delta
system might resolve this apparent paradox. The vegetation
change occurring in different parts of a fluvio-deltaic basin must
be taken into account for any future modeling effort.
10. Conclusions
(1) Sedimentology, chronology, isotope (d13C) and sequence
stratigraphic analysis of subsurface sediments of western delta
plain of Ganges–Brahmaputra delta shows that following the
deposition of marine clay of MIS 5, fluvial sands and mangrove
rich overbank muds (>23 ka age) deposited during the highstand of MIS 3.
(2) During the Last Glacial Maximum (LGM) sea-level lowering of
>100 m produced a regional unconformity (type 1), represented by palaeosol and incised valley. C4 vegetation expanded
on exposed lowstand surface in an ambient dry glacial climate.
The incised valley was filled by sands of 23–17 ka age during
the lowstand when the rate of sea-level fall decreased.
(3) At around w9 ka a rapid transgression inundated the lowstand
surface pushing the coastline and mangrove front w100 km
inland. Simultaneous intensification of monsoon and very high
sediment discharge (w4–8 times than modern) caused a rapid
aggradation of both floodplain and estuarine valley fill deposits
between 8 and 7 ka. The present Hoogli River possibly acted as
the main conduit for transgression and sediment discharge
that was subsequently abandoned. C3 vegetation dominated
the delta plains at this time.
(4) From 7 ka onward progradation of delta plain started and
continued till recent. This period experienced a mixed C3–C4
vegetation with localized mangroves in the mid-Holocene to
dominant return of C4 vegetation in the late Holocene period.
(5) The study indicates that while the initiation of the western part
of GB delta occurred at least 1 ka earlier than the global mean
delta formation age, the progradation started at w7 ka, at least
2 ka earlier than thought before.
(6) The terrestrial vegetation change was modulated by changes in
depositional environment, ecological niches and climate rather
than pCO2.
Acknowledgement
Cores for this study were raised under a joint collaborative project
between University College London and IIT, Kharagpur funded by the
Royal Society, U.K. Isotope data were generated in the mass spectrometer laboratory, IIT, Kharagpur funded by the DST, New Delhi. AS
thanks Prof. Rajiv Sinha, IIT, Kanpur for providing the grain size
analysis data of the samples and Prof. P. P. Chakraborty, ISM, Dhanbad
for his useful comments on an earlier version of manuscript. The
paper greatly benefited from suggestions of the editor Antony Long
and critical reviews of two anonymous reviewers.
Appendix 1. Supplementary data
Table S1. Summary of core locations and chronological data
generated and used in the study area; also given the available
chronology from the entire western delta plain. Table S2. Facies
types in the western GB delta plain. Table S3. Downcore data on
d13C, TOC (%), TN (%), C/N and d15N of bulk organic matter.
A. Sarkar et al. / Quaternary Science Reviews 28 (2009) 2564–2581
Appendix 1. Supplementary data
Supplementary data associated with this article can be found in
the online version at doi:10.1016/j.quascirev.2009.05.016.
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